Five samples from the IC and two samples from the EC were analysed. These were collected from internal (SE) to external areas (NW) of the Sesia Zone. A brief account is given here, with characteristic images shown in Fig. 2. Supplement S1 contains detailed descriptions and GPS locations.

The samples of the IC (FG1324, FG1315, FG12157, FG1347, and FG1249) are mica schists with eclogite facies assemblages comprising quartz, phengite, garnet, ± paragonite ± glaucophane ± omphacite ± chloritoid, with accessory allanite ± rutile. The main fabric in all of these samples is an eclogite facies foliation (Fig. 2a). Evidence of several deformation stages occurring before or after the main eclogite facies foliation is preserved in several samples as microlithons, commonly of phengite, omphacite, glaucophane, or chloritoid oriented at high angles relative to the main foliation, which wraps around them or is overgrown by them (Fig. 2b). Further evidence of several metamorphic stages occurring at eclogite facies conditions is reflected in growth zones of garnet (Giuntoli et al., 2018). Pre-Alpine relics include garnet cores (Fig. 2c) and zircon (cores ± first rims; Sect. 6.4 and 6.5).

Retrograde stages of blueschist or greenschist facies assemblages related to decompression are locally present in samples. The blueschist facies stage produced pleochroic crossite rims around glaucophane (Fig. 2d). The greenschist facies stage produced symplectites of actinolite ± albite ± chlorite around glaucophane and omphacite, chlorite at the expense of garnet, epidote, or clinozoisite rims around allanite, and titanite rims formed around rutile (Fig. 2e, f).

Sample FG1420, collected in the EC, is a garnet orthogneiss that shows a HP greenschist foliation marked by phengite, chlorite, and titanite; the foliation wraps garnet porphyroblasts that preserve a relic internal foliation (Fig. 2g). Permian magmatic relics of pleochroic allanite are surrounded by an Alpine corona of epidote grains (Fig. 2h). Some hundred metres to the north, another sample (FG12107) of the EC was collected. This is a leucogneiss characterized by the same metamorphic fabric and parageneses as the previous sample, except that garnet and magmatic allanite are missing.

Garnet and phengite display features in the IC samples that differ from those in the EC samples. To highlight and describe these, two main samples are compared in the following paragraphs: FG1249 (IC) and FG1420 (EC). A more complete account of garnet textures and mineral inclusions in the IC is shown in Giuntoli et al. (2018).

In the IC samples, garnet consists of a core followed by several rims with a grain size of up to several millimetres (Fig. 3a, b). The compositional map of the grossular endmember fraction (XGrs) in sample FG1249 shows a porphyroclastic core (Alm72Prp18Grs5Sps5; Table 1) with internal fractures sealed by garnet of higher XGrs (Fig. 3b). A first rim (Rim1-Alm76Prp15Grs9) overgrows the core and displays higher grossular contents. This Rim1 is followed by Rim2, which again records an increase in XGrs (Alm62Prp20Grs18). Rim2 resorbs parts of Rim1, externally and internally, and the core. Rim3 is peripheral and shows the highest Ca contents (Alm58Prp19Grs23).

Sample FG1315 is characterized by a porphyroclastic core (Alm69Prp28Grs4) with lobate edges and resorption features (details in Giuntoli et al. 2018 and Engi et al. 2018) surrounded by several rim generations: Rim1 (Alm61Prp21Grs19), Rim2 (Alm65Prp24Grs11), and Rim3 (Alm70Prp24Grs6). Atoll garnets, a few hundred microns in size, are observed in this sample. The shells of the atoll garnet have similar zoning patterns and compositions as the rim generations just described. In sample FG12157 the garnet core (Alm70Prp26Grs4) is rimmed by two growth zones: Rim1 (Alm64Prp20Grs16) and Rim2 (Alm59Prp24Grs17). In sample FG1347, the garnet core (Alm69Prp28Grs3) is enclosed by three rims (Rim1: Alm66Prp23Grs11; Rim2: Alm68Prp26Grs6; Rim3: Alm70Prp26Grs4). The exception is sample FG1324, in which garnet shows a single growth zone of homogeneous composition (Alm70Prp18Grs11Sps1).

In the EC, sample FG1420 shows garnet with completely different features. As shown in Fig. 3d, the XSps map highlights concentric zoning (values of Alm54Prp2Grs36Sps8 for the core, Alm57Prp3Grs37Sps4 for Rim1, Alm60Prp3Grs35Sps2 for Rim2, and Alm63Prp4Grs32Sps1 for Rim3), with no visible resorption features (further compositional endmember maps are shown in Supplement S2).

To link the growth zones of garnet to the main assemblage observed in the mineral matrix, microstructural relations, overprinting criteria, and mutual inclusions were employed, based on optical microscopy, SEM, and compositional maps. In particular, garnet in sample FG1249 contains inclusions of paragonite, phengite, and quartz between Rim1 and Rim2 (Fig. 3a). Rutile inclusions of a few microns are present in Rim2 and 3. Late chlorite fractures dissect the entire garnet. Garnet in sample FG1420 is wrapped by the main external foliation and includes an internal foliation marked by quartz, epidote, and titanite (Fig. 3c).

Phengite in IC samples displays a uniform composition except along grain boundaries, where lower Si and XMg contents are found, indicating retrograde overprinting (e.g. Fig. 3e; Group1 Si ∼ 3.36 apfu (atoms per formula unit), XMg ∼ 0.83; Group2 Si ∼ 3.3 apfu, XMg ∼ 0.68 in sample FG1249; Table 2).

In the EC, two distinct generations of phengite are distinguished based on their microtextural position: the first one describes the main foliation and is characterized by high Si values (Fig. 3f; Group1 Si ∼ 3.4 apfu, XMg ∼ 0.61 in sample FG1420), and the second phengite generation (Group2 Si ∼ 3.32 apfu, XMg ∼ 0.61) rims the first one and occurs in fold hinges that deform the main foliation.

For samples FG1324 and FG1420 the original bulk rock compositions obtained using XRF were used to compute isochemical phase diagrams. In samples FG1315, FG12157, FG1347, and FG1249, however, the unmodified bulk rock composition cannot be used for modelling because a significantly high volume fraction of garnet is present (5–10 vol %), including a pre-Alpine core and Alpine rim generations. To compute equilibrium diagrams properly, the reactive bulk rock composition was approximated using the program GrtMod (see Sect. 5.1.2). Each isochemical phase diagram is thus valid for a single P–T stage only. To select the reactive bulk composition of this specific stage, a link must be established between the particular garnet generation that formed in equilibrium with the mineral phases present in the matrix and the mineral phases. We established this link using petrographic observations, including textural equilibrium criteria, compositional zoning (visible in compositional maps), and inclusion relationships. Specifically, we determined that the garnet growth zones that coexisted with the mineral matrix are as follows: garnet Rim3 for sample FG1315, Rim1 for sample FG12157, Rim3 for sample FG1347, and Rim2 for sample FG1249. The corresponding reactive bulk compositions used for the modelling are provided in the Supplement S5.

Figure 4 shows isochemical P–T phase diagrams for each sample. The plots show the results of garnet thermobarometry (GrtMod), garnet isopleths (XGrs, XAlm, and XPrp in sample FG1324), and XMg and Si (apfu) isopleths for phengite. For sample FG1420, results of chlorite–white mica thermobarometry are also displayed. A summary of mineral compositional data for the main phases, the modelling method used, the XRF analyses of major elements of each sample, and details of the GrtMod results are available as Supplement S5–S7 and S8, respectively. Each sample from the IC and EC is presented separately below.

In the samples of the IC, allanite prisms are elongate in the eclogite foliation, showing mutual intergrowth relations with phengite, garnet, paragonite, and rutile, which define this main foliation (Fig. 7a). In samples FG1324, FG1315, and FG1347 allanite crystals are characterized by one main growth zone. Some thin (< 20–30 µm) allanite rims as well as epidote or clinozoisite rims that appear dark in BSE photos are observed (Fig. 7a, b); where present, these mark a retrograde greenschist overprint. In samples FG12157 and FG1249, BSE pictures show one or more allanite rims characterized by lower brightness (Fig. 7b, c). These rims may reflect minor retrograde stages that weakly altered the eclogite facies assemblage as well. Again, a peripheral epidote rim is present. Monazite is occasionally observed as a relic in allanite cores in samples FG1324, FG1315, FG1347, and FG1249. Monazite shows lobate edges and is surrounded by symplectites (micrometres in size) of allanite and apatite or by discrete crystals of apatite and allanite (Fig. 7d). These features suggest prograde growth of allanite and apatite at the expense of monazite, a common allanite-forming reaction (e.g. Janots et al., 2008).

Various mineral inclusions are found in allanite grains, as summarized in Table 5. In detail, allanite in sample FG1324 shows intergrowths with garnet (Fig. 7e), suggesting synchronous growth of the two minerals. Phengite inclusions analysed in allanite show the same chemical composition as those marking the main foliation (representative mineral analyses are available in Supplement S5). Based on these features and the alignment in the foliation, allanite is interpreted to have grown syn-kinematically in the foliation and at the same time as garnet. In sample FG1315 allanite includes phengite, paragonite, and garnet (Fig. 7a, f); the latter is similar in composition to the Alpine HP rims and atoll garnet. Phengite inclusions have the same composition as phengite marking the main foliation and as phengite included in atoll garnet. Due to the relationships of these mutual inclusions in this sample, allanite growth again appears to be related to the development of the main foliation. In samples FG1347 and FG1249 allanite includes phengite and paragonite; in FG12157 allanite includes only phengite. These micas have the same composition as those defining the eclogite foliation (Fig. 7b–d). In the cases of FG1249 and FG1347, allanite also shows intergrowths with both white micas (Fig. 7c).

In the EC, allanite is rare and, where present, is typically magmatic; it appears dark brown and pleochroic in the optical microscope, with a grain size of some millimetres (Fig. 2h; Giuntoli and Engi, 2016). In only two samples was metamorphic allanite found, preserved in the core of epidote crystals (FG1420, FG12107). The metamorphic allanite has a grain size of less than 50 µm and is colourless or pale yellow in polarized light, with low interference colour and undulose extinction in crossed polarized light. Sample FG1420 shows both magmatic and metamorphic allanites (Fig. 7g, h). Allanite includes paragonite, with phengite and titanite occurring both in the epidote rim and at the allanite–epidote boundary (Fig. 7g). Very few tiny monazite grains (a few micrometres) are found in the core of metamorphic allanite. Sample FG12107 also shows similar epidote crystals preserving metamorphic allanite in their core, as in sample FG1420. Magmatic allanite preserved in sample FG1420 occurs as mm-size grains that are fractured and appear much brighter in BSE pictures than metamorphic allanite (Figs. 2h, 7h). Epidote crystals form satellites around magmatic allanite, suggesting partial breakdown (Fig. 7h). Note that these epidote crystals retain a BSE bright core of newly grown (Alpine) metamorphic allanite.

For both IC and EC, only allanite cores were successfully dated. The rims showed too high common lead (Pbc) contents, the correction (Gregory et al., 2007; Burn et al., 2017) of which would lead to large uncertainties in the age calculation.

In the IC and EC, the Tera–Wasserburg and 232Th / 206Pbc-208Pb / 206Pbc isochron diagrams are concordant, within the range of uncertainty, in all the analysed samples (Figs. 8, 9; every ellipse is a single spot measurement). Ages range between 77 and 56 Ma for the allanite cores of the IC. In the EC, the magmatic allanite cores in sample FG1420 yield ages of ∼ 290 Ma, and the metamorphic cores yield ages of 73.7 ± 8.2 Ma. In sample FG12107, the metamorphic allanite cores were dated to 62.8 ± 3.3 Ma. A summary of the Alpine allanite age data from each sample is listed in Table 5.

In detail, the IC Tera–Wasserburg diagrams show 207Pb / 206Pb y intercepts of 0.84 ± 0.01, 0.823 ± 0.018, and 0.83 ± 0.004 for samples FG1324, FG1315, and FG12157, respectively, with a MSWD on the regression between 1.2 and 2.1 (Fig. 8). 206Pbc–isochron diagrams display a 208Pb / 206Pb y intercept of 2.081 ± 0.027, 2.077 ± 0.066, and 2.085 ± 0.028 for samples FG1324, FG1315, and FG12157, respectively, with a MSWD on the regression between 0.46 and 0.98. These values are close to the predicted values of Stacey and Kramers (1975) for model lead evolution of this time range (Fig. 1 in Burn, 2016). The exception is sample FG1347, in which the Tera–Wasserburg diagram shows a 207Pb / 206Pb y intercept at 0.787 ± 0.04 (MSWD on the regression of 2.5) and displays a 208Pb / 206Pb y intercept on the 206Pbc–isochron diagram at 1.98 ± 0.082 (MSWD on the regression of 0.4). These values deviate from the values predicted by Stacey and Kramers (1975), probably indicating local inheritance.

In the EC, data for magmatic allanite in sample FG1420 show a 207Pb / 206Pb y intercept at 0.85 ± 0.11 in a Tera–Wasserburg diagram and display a 208Pb / 206Pb y intercept at 2.09 ± 2.7 in a 206Pbc–isochron diagram (Fig. 9). These large uncertainties reflect few (eight) spots analyses. Metamorphic allanites show 207Pb / 206Pb y intercepts of 0.825 ± 0.01 and 0.824 ± 0.006 in Tera–Wasserburg diagrams (MSWD on the regression of 0.7 and 1.2) and display 208Pb / 206Pb y intercepts of 2.061 ± 0.029 and 2.089 ± 0.022 in 206Pbc–isochron diagrams (MSWD on the regression of 0.4) for samples FG1420 and FG12107, respectively. These values are close to the predicted values of Stacey and Kramers (1975) for these time ranges.

Internal textures of zircon from the IC reveal complex zoning in CL images (Fig. 10), showing detrital cores and several phases of resorption and (metamorphic) overgrowth. Zircon cores commonly preserve a variety of internal textures, most commonly oscillatory zoning (Fig. 10e, f), which is typical of zircon-grown form melt. Many cores are affected by resorption, obliterating earlier features, but some grains show sharp boundaries between core and rims, and these preserve features of sediment transport such as broken, rounded, or pitted surfaces (Fig. 10b). The number of rims varies among and within samples, from two rims in sample FG1257 to a maximum of five in FG1315. Most zircon grains show a first metamorphic rim with a light grey to bright CL response, followed by a rim with dark CL appearance. The third rim is again typically medium grey to bright white in CL. In sample FG1315, a fourth (dark CL) and fifth (light grey) rim follow, and FG1347 occasionally shows a CL-dark fourth rim. The internal textures of the different rims are not always clearly distinguishable, either because of limited width or, in the case of a very dark or bright CL response, limited contrast. The second metamorphic rim in FG12157 is either uniform or shows fir-tree zoning (Vavra et al., 1996; Root et al., 2004). In sample FG1347 metamorphic rims are often too thin to distinguish or no texture is recognizable in the bright CL third rim. The first two rims in sample FG1315 have cloudy textures; the third rim with the bright-CL shows no further internal textures but sometimes has inclusions and the two outermost rims commonly are uniform or cloudy in texture. The innermost and outermost metamorphic rims in samples FG1249 and FG1324 are bright and featureless; the second dark rim in sample FG1249 shows sector zoning and many inclusions. The third (medium grey) rim in FG1249 has a cloudy texture.

Zircon in the EC displays pre-Alpine magmatic oscillatory zoning, rarely shows a bright CL rim, is 5–10 µm in width, and surrounds the core.

The range of Alpine 206Pb / 238U zircon dates for each sample are summarized in Table 5. Details on the pre-Alpine dates are available in Kunz et al. (2017). Supplement S10 gives an overview of pre-Alpine dates as well as individual analyses of Alpine dates. Concordia plots of individual spot analyses of Alpine zircon are available in Supplement S11. For the five samples we obtained 27 analyses of detrital cores, 53 of Permian rims, and 34 of Alpine rims. Detrital zircon cores in samples from the IC range from ∼ 793 to 353 Ma and partially overlap with the first rim in samples FG1324 and FG1249 (Fig. 10), for which the dates range from ∼ 450 to 420 Ma. In most samples, at least two rims are found yielding Carboniferous to Triassic dates from ∼ 313 to 222 Ma, as discussed in detail by Kunz et al. (2017); only one sample (FG1324) shows none of these. Alpine rims, with a range between 77 and 56 Ma, were measured in all samples except FG1249, in which the third rim was too thin (10 µm) to be analysed. No correlation was found between rim types and ages within samples or amongst them, and it was not possible to make absolute age distinctions among Alpine rim generations. This is due to relatively large uncertainty in the ages, as the narrow rims only allowed small spots for LA-ICP-MS measurements, decreasing their analytical precision.

Th / U ratios in detrital zircon cores range between 0.15 and > 3; the rim generation with dates between 450 and 420 Ma have low Th / U ratios in sample FG1324 (0.006–0.14) whereas those in sample FG1249 are between 0.1 and 0.5. The Carboniferous to Triassic rims show Th / U ratios between 0.002 and 0.4. The Alpine metamorphic rims show low Th / U ratios between 0.001 and 0.01.

Preliminary analysis of allanite for U–Th–Pb using LA-ICP-MS was performed in situ (i.e. using LA-ICP-MS spot analyses) on polished thin sections. However, to obtain more material for dating, allanite grains were subsequently separated, mounted on resin mounts, and polished. The main reasons why allanite grains were separated for dating are as follows.

Grain mounts are more efficient as they enable the selection of more spots to analyse a single growth zone (at least 1 order of magnitude more spot analyses per sample than in thin sections).

The fundamental task to reconstruct the P–T–t paths is to establish a strong link between P and T conditions, derived from thermodynamic modelling, and t data, derived from age dating. Several criteria were used to achieve this goal, notably (i) textural evidence, (ii) the presence and composition of distinctive minerals as inclusions, and (iii) the presence of intergrowths of allanite with other phases.

As described in Sect. 6.2, allanite grains of the IC are intergrown with the main mineral phases that describe the eclogite facies foliation (Fig. 7; Table 5). Furthermore, the compositions of mineral inclusions in allanite matched those found in the matrix and those predicted by thermodynamic modelling (representative chemical analyses in Supplement S5). Based on these observations (and details given in Sect. 6.2 for each sample), we infer growth of allanite coeval with the minerals marking the eclogite facies foliation in all of the samples analysed from the IC. This link between age data and P–T conditions is shown by the red ellipses on the equilibrium phase diagrams shown in Fig. 4 and on the P–T–t paths shown in Fig. 11b and c (colour-coded for each sample). Additionally, in four samples of the IC, the ages obtained for allanite cores and Alpine zircon rims overlap within uncertainty, suggesting coeval growth of these two accessory minerals.

In the EC, metamorphic allanite is rare and preserved only in the core of epidote (Sect. 6.2; Fig. 7g, h). Epidote is intergrown with white mica, chlorite, and titanite, marking the retrograde greenschist overprint, which is constrained by chlorite + white mica + quartz + H2O thermobarometry (∼ 0.9 GPa and 350 ∘C to 0.4 GPa and 230 ∘C; see Fig. 6). According to these results, metamorphic allanite growth in the EC predates the retrograde greenschist overprint (red ellipse in Fig. 4f and purple ellipse in Fig. 11b–c).

The IC shows several tectonic sheets, from several hundred metres to a few kilometres in thickness (Giuntoli and Engi, 2016), some of which moved independently in the subduction channel (Rubatto et al., 2011; Regis et al., 2014) at some stages of the evolution. Some of the samples we studied, while taken at most a few kilometres apart in the field, recorded similar P–T paths at different times, as reflected by the three age groups identified (Fig. 1b, c). This age difference may reflect relative tectonic mobility at eclogite facies conditions between such sheets of basement-derived rocks, which are notoriously difficult to delimit in this terrain (Giuntoli and Engi, 2016). In addition to tectonic mobility, the combined effects of deformation and repeated hydration must be considered as triggers of recrystallization that promoted chemical equilibration to eclogite facies assemblages (including allanite and zircon, as discussed below).

The differences in the P–T-conditions (at eclogite facies) documented above for the IC samples do not indicate trajectories as tortuous as those suggested by numerical models for ablative subduction (e.g. Stöckhert and Gerya, 2005; Roda et al., 2012). Our P–T–t data as well as the nature and geometry of tectonic boundaries mapped by Giuntoli and Engi (2016) rather indicate progressive accretion at depths of 50–60 km, probably during final subduction–early exhumation stages, juxtaposing a series of continental sheets and forming a coherent complex (the IC; Vitale Brovarone et al., 2013; Regis et al., 2014). Eclogite facies conditions in the IC evidently prevailed for an extended period, at least ∼ 15 Myr.

In the IC, the main deformation stages and mineral reactions were related to pulses of external fluid passing through the rocks, as reported by Engi et al. (2018) and Giuntoli et al. (2018). Repeated fluid influx occurred under eclogite facies conditions, as shown by resorption and growth features in garnet and zircon, with hydrous fluid pervasively rehydrating rocks that had previously been almost completely dehydrated by upper Paleozoic metamorphism (Engi et al., 2018). Thus Permian kinzigites (granulite facies metapelites) were transformed back to mica schists during Alpine subduction. As allanite and zircon ages from these samples are identical within analytical uncertainties, it appears that this fluid influx also triggered the coeval growth of accessory phases under eclogite facies conditions. Later and more localized fluid influx has also been documented in blueschist facies conditions (Konrad-Schmolke et al., 2011), and this probably continued to greenschist facies conditions, as reflected by the partial overprint of the main eclogite assemblages, which is frequently observed in the IC. Our samples in the IC indicate higher temperatures than have been previously reported (Fig. 11). This discrepancy may reflect a regional thermal gradient, namely higher temperatures in more external areas (NW) of the IC, but we suspect it to be a direct effect of the (re)hydration of pre-Alpine HT rocks, which is an exothermic process (e.g. Walther and Orville, 1982; Peacock, 1987; Lyubetskaya and Ague, 2009). Walther and Orville (1982) analysed the thermal effects of (de)hydration reactions during regional metamorphism and found them to be substantial. When applied to the present situation of (re)hydration, heat capacity data indicate that heating the HT Permian protolith requires ∼ 1.0 kJ K-1 kg-1 of leucogneiss. The enthalpy released upon partial hydration of this protolith (producing the water content typical of these mica schists, 1.5 wt % H2O) adds ∼ 77 kJ kg-1 in enthalpy. Such hydration should thus result in a temperature increase of some 80∘. This estimate lends credibility to the P–T estimates obtained here, which are indeed 20–80 ∘C higher than some maximum temperatures recently reported from other parts of the Sesia Zone, e.g. 575 ∘C (Konrad-Schmolke and Halama, 2014), 570–630 ∘C for the Druer Slice (Regis et al., 2014), or > 600 ∘C for the Ivozio Complex (Zucali and Spalla, 2011). In addition, Zr-in-rutile data reported by Kunz et al. (2017, their Table 3) gave 640 ∘C for one of the present samples (FG1249), confirming our results by an entirely independent method.

A counterclockwise P–T path is proposed here for the EC. This trajectory is based on the results of garnet and chlorite–white mica modelling. The path as shown in Fig. 11 implies that, under nearly isobaric conditions, this area of the EC experienced cooling by 100–150∘, possibly related to the entry of cold material into the subduction channel, as already proposed by Pognante (1989) for the southern Sesia Zone. Piemonte–Liguria oceanic units are a likely source of such material, which would be in line with age data of 58–40 Ma for the HP metamorphism in the Zermatt–Saas Zone (e.g. Rubatto et al., 1998; de Meyer et al., 2014; Weber et al., 2015) and with tectonic–kinematic models for the evolution of the Sesia–Dent Blanche nappes (Angiboust et al., 2014; Manzotti et al., 2014).

Summarizing, the major differences between the eclogite facies conditions recorded in the IC and the epidote blueschist facies conditions in the EC are now quantitatively constrained by P–T–t data presented in this study (Fig. 11). Tectonic juxtaposition of these two complexes appears to have occurred during exhumation probably at 0.7–0.9 GPa and 350–400 ∘C since these are the first joint P–T conditions found in both complexes around 46–38 Ma (age data from Inger et al., 1996, and Reddy et al., 1999). These data confirm the proposition by Williams and Compagnoni (1983) of a first-order tectonic contact (Barmet Shear Zone in Giuntoli and Engi, 2016; Fig. 1b, c) between the two complexes. Our data are the first to quantify the P–T discontinuity at that contact, i.e. ∼ 1 GPa and 100–180 ∘C. This age range is younger compared to the age range proposed in the tectonic model by Angiboust et al. (2014) for the juxtaposition of the two complexes (70–57 Ma, based on data of Babist et al., 2006; Konrad-Schmolke et al., 2011; and Halama et al., 2014). Whereas Angiboust et al. (2014) considered the Tallorno–Chiusella Shear Zone (Babist et al., 2006; Konrad-Schmolke et al., 2011) to represent the contact between the IC and EC, more recent mapping (Giuntoli and Engi, 2016) shows that this contact is located several kilometres further NW, in the Barmet Shear Zone.

The data shown in Fig. 11 allow us to derive average exhumation rates for the IC (Table 6). Using Groups 1 and 3 as the two anchors, i.e. the oldest and the youngest groups, a first calculation considers Stage1, the interval from the highest pressures recorded in these two groups (∼ 1.6 GPa at ∼ 73 Ma and ∼ 2 GPa at ∼ 55 Ma) to the greenschist conditions upon juxtaposition with the EC (∼ 0.6 GPa at ∼ 38 Ma, age from Inger et al., 1996). The mean vertical exhumation velocity obtained ranges from 0.9 to 2.7 mm year-1 (Table 6). If we take the juxtaposition to be at ∼ 46 Ma (Reddy et al., 1999) for the same P conditions, values for the mean vertical exhumation velocity increase to 1.2 and 5.1 mm year-1. Plate convergence between Adria and Europe was ∼ 15 mm year-1 in the time span of 68–38 Ma (Handy et al., 2010). In the above calculations we considered subduction angles of 90, 60, and 45∘ and a thickness of 20 km for the upper crust and 10 km for the lower crust; their densities were taken as 2.7 and 3.0 g cm-3, respectively, and 3.3 g cm-3 for the upper mantle.

For Stage2, the final stage of exhumation of the IC and EC, we obtained mean vertical exhumation velocities using the juxtaposition conditions (∼ 0.6 GPa) and ages quoted for the end of Stage1 and 32.5 Ma for the arrival at the surface. That age of the Biella Volcanic Suite (Kapferer, 2010; Kapferer et al., 2012) is supported by zircon fission track ages (Berger et al., 2012). Depending on the time of juxtaposition adopted, at 38 or 46 Ma (Table 6), a mean vertical exhumation velocity of 4 or 1.6 mm year-1 results. For Stage2, only the vertical exhumation velocity was computed, and an average density of 2.7 g cm-3 was assumed. The rate of convergence in the 35–20 Ma period was ∼ 13 mm year-1 (Handy et al., 2010). Data for Stage2 are taken as an approximate range of exhumation rates since the last part of the exhumation was not uniform throughout the Sesia Zone; owing to brittle structures local differences may be important (e.g. Berger et al., 2012; Malusà et al., 2006). Furthermore, the values given are minimum exhumation velocities, as the emplacement of the Biella Volcanic Suite may have post-dated the arrival of the Sesia Zone at the surface.

The Stage1 exhumation rates we obtained for the IC agree with exhumation velocities proposed for the same complex by Zucali et al. (2002) and Regis et al. (2014) and the mean exhumation rates for the Sesia Zone proposed by Roda et al. (2012). However, our rates are up to an order of magnitude lower than those proposed by Rubatto et al. (2011) and the maximum values of Roda et al. (2012). Mean exhumation velocities for the final exhumation stage of the Sesia Zone (Stage2) are in the same range (1.6–4 mm year-1) as those for Stage1 (0.9–5.1 mm year-1). Our data show no decrease in exhumation velocity from early exhumation (i.e. mantle to deep crustal positions) to late stages (i.e. crustal levels to the surface), such as has been proposed by Rubatto and Hermann (2001) for ultra-HP terranes.

Our work highlights that subducted continental terranes can be composed of several complexes that experienced major differences in their subduction histories (i.e. ∼ 1 GPa in pressure and 100–180 ∘C in temperature for the Sesia Zone). These complexes could have been tectonically juxtaposed during late metamorphic stages (i.e. greenschist facies conditions) before being jointly exhumed to the surface. Complexes can be lithologically heterogeneous and may comprise several tectono-metamorphic subunits (from a few hundred metres to several kilometres in thickness). These may experience similar P–T paths but at different times (5–10 Myr apart). Differences in the P–T–t trajectories would thus reflect different tectonic sheets and attest to tectonic mobility in the subduction zone and/or several stages of internal deformation plus hydration at eclogite facies that triggered a pervasive HP fabric and assemblage (including datable accessory minerals).

To unravel such complex histories in subducted continental terranes, carefully established field relations, followed by microstructural and petrochronological analysis need to be combined to

map and identify the major and secondary tectono-metamorphic contacts;

Finally, the heterogeneity of subducted continental terranes highlighted by this study should be considered when comparing results to numerical and analogue models that aim to investigate the mechanisms responsible for the subduction and exhumation of the continental crust.